DISTRIBUTION AND EVOLUTION OF 'LATERITES' AND LATERITIC WEATHERING
PROFILES, DARLING RANGE, WESTERN AUSTRALIA
R. R. Anand and C.R. M. Butt
CRC for Landscape Environments and Mineral Exploration, CSIRO Exploration and Mining, 26 Dick Perry Avenue,
Technology Park, Kensington, Western Australia 6151.
ABSTRACT
This paper summarizes the distribution and characteristics of the lateritic regolith of the Darling Range and presents
models for its formation and evolution. Typical, complete, weathering profiles on granite average about 20 m in
thickness and consist of gravelly soil, lateritic duricrust, saprolite and saprock. Lateritic duricrusts occupy gently
sloping to horizontal upland areas and are either residual, or locally transported and recemented. In much of the Darling
Range, the lower part of the duricrust, especially on hill slopes rather than crests or valley floors, is highly aluminous
and forms an extensive resource of bauxite. Fragmental, fragmental-pisolitic, pisolitic and vesicular types can be
identified on the basis of secondary structures. Fragmental duricrust largely consists of gibbsite, hematite, goethite and
quartz and has resulted from direct gibbsitization of saprock or bedrock without forming the kaolinite-rich deep
saprolite. Outcrops of duricrust with relict bedrock textures are common. In contrast, pisolitic duricrust with hematitemaghemite and χ-alumina rich mineralogy have a more complex history than fragmental duricrust with simple
mineralogy. Vesicular duricrust is goethite-rich and is formed by the ferruginization of sandy detritus and quartz
pebbles. The profiles show no condensed sequences, and individual rock types are traceable geochemically and
mineralogically, but with increasing difficulty, from bedrock to the surface. The concentrations of Fe, Al, Si, Ti, V, Cr
and residual quartz, particularly in fragmental duricrust, can be used to identify bedrock. Deep weathering profiles at
Jarrahdale and Boddington yield late Tertiary palaeomagnetic ages and it appears that modification of these profiles to
form bauxite is continuing today.
1
INTRODUCTION
The Darling Range is a dissected, uplifted plateau developed on the western margin of the Archaean Yilgarn Craton and
some adjacent Mesozoic rocks of the Perth Basin. It extends from approximately the Moore river 100 km north of
Perth to Collie and Bridgetown, 220 m south. To the west, it is bounded by an escarpment, which approximately
follows the Darling Fault, separating the Yilgarn Craton and the coastal plain. The eastern boundary is less clearly
defined, but in part follows the lowlands, including the Avon valley, marking the Cape Riche-Yandanooka lineament. A
deeply weathered lateritic regolith is developed throughout the range, on igneous and metamorphic Archaean rocks and
sediments of Permian and Mesozoic age. The regolith, which varies in thickness from a few metres to over 50 m, is
generally well preserved, except in the valleys of the major drainages that pass through the range (Swan-Avon, Murray,
Collie and Blackwood rivers). It is also being eroded by smaller streams and rivers that rise in the Range itself. From
east of Perth to the Harris River area, north of Collie, the regolith is bauxitic and forms the world's leading alumina
producing region.
Lateritic duricrusts are characteristic and important features of the Darling Range. They occupy gently sloping to
horizontal uplands and are considered to have formed by in situ weathering of parent rocks (Baker, 1972; Sadleir and
Gilkes, 1976; Davy, 1979; Murray, 1979; Butt, 1981; Davy and El-Ansary, 1986; Anand et al., 1991; Hickman et al.,
1992; Anand, 1994; Anand and Paine, 2002) and weathering of fluvial (Grubb, 1971) or aeolian sediments (Brimhall et
al., 1988; Glassford and Semenuik, 1995). There have many discussions on the role of parent rock, climate,
topography, tectonic history and vegetation in the development of bauxitic profiles (Terrill, 1956; Playford, 1959;
Grubb, 1971; Baker, 1972, 1976; Sadleir and Gilkes, 1976; Davy, 1979; Hickman et al., 1992). The objectives of this
paper are to: (i) summarize the distribution and characteristics of regolith materials in the Darling Range including
lateritic duricrusts and (ii) present models of their development and evolution.
2
CLIMATE AND VEGETATION
The Darling Range has a typical, highly seasonal, Mediterranean climate, with cool, wet winters and warm dry
summers. The mean maximum temperature ranges between 12 and 28°C. The total rainfall is closely related to
latitude, distance from the coast and local relief and elevation. Rainfall ranges from over 1300 mm per annum on high
ground (>300 m AHD) near the scarp to below 460 mm per annum in the Avon valley near York. The grade of bauxite
decreases progressively with the present day rainfall of the area so that, where rainfall is less than 600 mm, bauxite
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grade is below the level of commercial exploitation. The dominant vegetation is open eucalypt forest. In wetter areas,
near the scarp, varying proportions of Eucalyptus marginata (jarrah) and E. calophylla (marri) form the upper storey on
the ridge crest and slopes. In drier areas inland, the woodlands are more open and characterized by abundant E.
wandoo and E. accedens (wandoo). Banksia grandis, Persoonia longifolia, Xanthorrhoea preisii and Macrozomia
riedlei dominate the middle and lower storey. E. calophylla and Melaleuca preissianan dominate the upper storey of
the valley floors, where the lower storey contains Xanthorrhoea preisii, Kingia australis and Hypocalymma
angustifolium.
Figure 1. Geomorphological divisions of the Darling Range and adjacent areas (Hickman et al., 1992).
3
BEDROCK GEOLOGY
Much of the Darling Range is underlain by crystalline rocks of Western Gneiss Terrane of the Archaean Yilgarn Craton.
The rocks are principally granites and granitic gneisses that have been intruded by numerous northwest trending late
Proterozoic tholeiitic, quartz dolerite dykes that commonly constitute over 10% of bedrock (Hickman et al., 1992). The
oldest units (>3.0 Ga) are the Jimperding Gneiss Complex and Chittering Metamorphic Belt in the north, and the
Balingup Gneiss Complex in the south. These consist dominantly of orthogneiss and paragneiss, with metamorphic
assemblages of amphibolite to granulite grade. The remaining granitic terrain is younger, 2.55-2.7 Ga, and of low
metamorphic grade, and has intruded mostly older (2.65 Ga) greenstone rocks of the Saddleback Group. The
Saddleback Group extends for 43 km north-northwest from near Boddington and is between 5 km and 12 km wide. It
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consists of a weakly metamorphosed felsic and mafic volcanic and pyroclastic rocks, with minor sedimentary units,
largely fault-bounded against granite and migmatite.
Post-Proterozoic sedimentary rocks in the Darling Range include coal-bearing Permian sandstones and siltstones that
occupy small basins in the Archaean rocks at Collie and Wilga, and up-faulted Mesozoic sediments of the Perth Basin
that form the Dandaragan Plateau, north of Perth, between the Gingin scarp, which follows a western splay of the
Darling Fault, and the Darling Fault itself. There are also small but widespread occurrences of deep unconsolidated
quartzose sediments, sometimes containing rounded siliceous gravels (up to 7 m thick). These may occur along shallow
valleys or valley divides as well as on low crests (Mulcahy et al., 1972).
4
GEOMORPHOLOGY
The Darling Range has been divided into three geomorphological provinces (Churchward and McArthur, 1980;
Hickman et al., 1992), namely the Dandaragan Plateau in the north. the Darling Plateau and the Blackwood Plateau in
the south (Figure 1). This discussion is confined to the Darling Plateau.
Figure 2: Geomorphological features of the Darling Range. (A) Intensity of slope about the Boddington district with
slopes ranging from 0° to 35°. (B) Shaded elevation model about the Boddington district (data courtesy of the
Australian Surveying and Land Information Group, Canberra, Australia. Crown Copyright©. All rights reserved).
The western boundary of the Darling Plateau, the Darling Scarp, rises abruptly 100 to 300 m above the low-lying
subdued topography of the Swan Coastal Plain. The position of the scarp closely corresponds to that of the Darling
Fault, which separates resistant Precambrian rocks to the east from Phanerozoic sedimentary rocks of the Perth Basin.
The Darling Plateau has an undulating surface (e.g. Boddington district, Figure 2A, B), with crests ranging in elevation
from 280 to 330 m and swampy valley floors set some 50 to 100 m below. Projecting above the general plateau level
are summits rising to over 500 m, the most prominent of which are Mt. Cooke (582 m), Mt. Dale (543 m), Mt. Randall
(523 m), Mt. Solus (576 m), Wourahming Hill (503 m) and Mt. Saddleback (575 m). The terrain is dissected in the east
and more so in the west, a trend that relates to progressive drainage incision from shallow swampy tracts on the plateau
to deep valleys near the Darling Scarp. Dissection of the plateau has created three distinct subdivisions: lateritic
uplands, deeply incised valleys and wide valleys with intervening low hills (Figure 1).
The lateritic uplands are mantled with deeply weathered profiles capped with lateritic duricrust, lateritic gravel and
sand and form a undulating surface; sand generally occurs in shallow depressions at the heads of drainages. Deeply
incised valleys are restricted to the western part of the plateau where streams and rivers have cut narrow, steep-sided
valleys through the surface of the lateritic upland. Erosion is commonly sufficiently deep to expose bedrock in the
lower slopes and valley floors. Wide valleys and intervening areas of low hills occur in the eastern part of the Darling
Plateau. Small areas of poorly-cemented lateritic duricrust occur on interfluves; bedrock is locally exposed in the more
hilly areas (Hickman et al., 1992). Topography is generally subdued although deep valleys are associated with major
drainages such as the Avon, Murray and Collie rivers and their larger tributaries.
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ANAND & BUTT
WEATHERING PROFILES
The interplay of weathering, erosion and relocation of materials, generally over long periods, has formed the lateritic
regolith that mantles the Darling Range. On granite, the regolith (Figure 3) commonly averages about 20 m in thickness
and consists of gravelly soil, lateritic duricrust (with a lower, bauxitic, zone in wetter areas), saprolite and saprock.
Similar profiles form on greenstones, but are generally deeper (>40 m). However, in detail, profiles are extremely
variable – for example, considerable thicknesses of kaolinitic clay form where porosity is greater over shears and faults.
In places, transported duricrust overlies fresh bedrock.
Soil forms the upper part of the regolith. It is generally 0.2-2.0 m thick (mean ~0.5 m) and consists of lateritic gravel in
a clayey sand matrix. The soil overlies lateritic duricrust, which reaches 1-3 m in thickness over granite and up to 4 m
over mafic rocks. Over granitoids, the duricrust is light-brown, with abundant visible quartz; over dolerite dykes, it is
red-brown and contains very little quartz. Pisolitic duricrust forms the upper unit of the duricrust and consists of black
to red-brown, simple or compound, nodules and pisoliths cemented by gibbsite and goethite. This unit may be either
essentially residual, or locally transported and cemented. The underlying fragmental duricrust is commonly residual, as
indicated by clasts that preserve primary lithic fabrics, such as the original quartz grains and feldspar lathes
pseudomorphed by gibbsite of granitoid bedrock. The friable bauxite zone immediately underlies the fragmental
duricrust. This zone is yellowish to reddish brown, generally about 3 m thick but can be up to 10 m thick in mid slope
positions on large ridges. The bauxite zone contains round to angular nodules and saprolitic fragments in a fine grained,
loose, earthy or sandy matrix. The coarser material ranges from 2 mm to over 100 mm. Feldpars pseudomorphs are
common in gibbsite nodules and are generally 1-2 mm in size.
The contact between the bauxite zone and saprolite is commonly abrupt. The saprolite is generally 10-15 m thick over
granitoids, generally mottled by Fe oxides in the upper part, but becoming increasingly bleached with depth. It consists
of vermicular kaolinite pseudomorphs after primary grains of mica, an isotropic groundmass of kaolinite after feldspar,
coarse quartz and abundant large voids. The saprolite thus has a much lower bulk density and higher porosity than the
parent rock. Over mafic rocks, saprolite ranges from white to multicoloured kaolinitic clay with mottles, commonly
overprinted by liesegang rings. Saprolite passes downwards into saprock, before fresh bedrock is reached. The
weathering front on dolerite commonly occurs over a few centimetres, whereas over granitoids, it is generally
gradational over 2 - 3 m. On a local scale, the topography of the weathering front is much more irregular than that of
the duricrust. Pinnacles of bedrock and isolated corestones may occur high in the weathered profile.
On steep slopes, the lateritic profile may be partly or completely eroded, or have never fully formed. In such situations,
there may be a surficial transported unit, commonly containing lateritic gravels derived from upslope, overlying
weathered or fresh bedrock. The soil is developed from the transported material or, in its absence, directly from
saprolite or bedrock.
The economically significant bauxite zone is not always present even where the regolith is fully preserved. It is best
developed in freely drained slopes, becoming less aluminous (i.e. not bauxitic) on crests, lower slopes and valleys. As
noted above, it also has formed only in higher rainfall areas of the Range between Perth and Collie. Farther south, east
and north, it is less abundant. However, an exception to the overall pattern of bauxite becoming less common eastwards
is the anomalously high bauxite concentration developed on the Mount Saddleback greenstone belt, where bedrock
composition and topographic conditions have combined to produce locally favourable conditions for bauxite formation.
6
MINERALOGY OF WEATHERING PROFILES
The mineralogical compositions of bauxitic laterite profiles developed on granite, andesite and metabasalt are compared
in Figure 4. Iron oxides and gibbsite are the dominant minerals in the lateritic duricrust and bauxite zone, whereas
kaolinite and/or halloysite dominate the saprolite. Quartz is present throughout the regolith on granitoids. Goethite,
hematite and maghemite abundances generally increase towards the top of the profile. They may be very
heterogeneously distributed, being absent from certain horizons, yet being dominant minerals only a few centimetres
below. They are more abundant on mafic than in granitic profiles. Goethite may occur anywhere in the profile, but is
especially common in the upper horizons. On granite, the abundance of goethite generally decreases sharply at the
lower contact and is absent from the underlying white saprolite. Where the profile is derived from mafic rocks, as at
Boddington, goethite persists though the saprolite. Hematite, although much less abundant is nevertheless common,
particulary in the duricrust. Maghemite is present in near-surface pisolitic duricrust and loose pisoliths, commonly
with traces of corundum, associated with relatively high concentrations of hematite and low concentrations of kaolinite.
The maghemite-rich pisoliths contain large amounts of Al2O3 (40-50%), which indicates that maghemite probably did
not form from solution, since this is inhibited by dissolved Al (Taylor and Schwertmann, 1974; Coventry et al., 1983.
Maghemite and corundum can be formed by dehydroxylation of Fe and Al oxyhydroxide by heating to approximately
300-500°C in the presence of organic matter (Anand and Gilkes, 1987a). Field relationships, including the association
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of hematite, maghemite and corundum, suggest this to be the probable mechanism, with suitable conditions provided by
ancient forest fires (Schwertmann and Fechter, 1984; Milnes et al., 1985; Anand and Gilkes, 1987a)
Figure 3: Typical lateritic bauxite weathering profile. Photomicrographs of fabrics from various horizons of a profile
developed over granite at Jarrahdale. (A) Optical micrograph of a polished section of pisolitic duricrust showing
hematite-maghemite-rich (1) and hematite-rich pisoliths (2) surrounded by lighter (3) and darker (4) cutans in a quartzgibbsite-rich matrix (5). (B) Fragmental duricrust showing angular to sub angular hematite-gibbsite-rich fragments (1)
in a gibbsite-quartz-rich matrix (2). (C) Optical micrograph of upper saprolite showing feldspars completely replaced
by kaolinite (1) set in a close packed relict quartz (2). (D) Optical micrograph of middle saprolite showing the
alteration of feldspar to a mixture of halloysite and kaolinite (1), highly altered biotite which has exfoliated and goethite
and/or clay minerals between the cleavage planes (2) and opaque mineral (3). (E) Optical micrograph of saprock
showing early stages of alteration of feldspar to a mixture of halloysite and kaolinite (1) and partly altered biotite
showing a mixture of Fe oxide and clay minerals (2) and quartz (3).
Gibbsite occurs in significant amounts in the bauxite zone and the fragmental duricrust, but decreases appreciably in
pisolitic duricrust and pisolitic lag gravels where poorly crystalline alumina (PCA) becomes the major Al mineral
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(Figure 5). Gibbsite consists mostly of 0.1-2 μm platy crystals (Anand et al., 1991); some occurs as subhedral to
euhedral hexagonal crystals, but most is present as components of basally oriented aggregates. The PCA appears to be a
major constituent in near-surface pisoliths and nodules and decreases markedly or is absent in subsurface fragmental
duricrust (Anand, 1994). The trend in PCA distribution in the profile is similar to that of hematite and maghemite, but
is opposite to goethite and gibbsite. PCA has low loss of ignition (LOI), which is thus an indirect indication of the total
gibbsite, goethite and kaolinite content. Two distinct populations for Al2O3 versus LOI, one with a positive correlation
and one with a negative correlation (Figure 6) suggest that, for a given abundance of Al 2O3, the LOI is greater in
fragmental duricrust and bauxite than pisolitic duricrust and lag, suggesting that most Al in the latter is present as PCA,
not gibbsite. Singh and Gilkes (1995) subsequently demonstrated that this unaccounted alumina is χ-alumina, one of
several poorly crystalline forms of anhydrous alumina.
Boehmite appears to occur only in small amounts in loose pisoliths and pisolitic duricrust (Grubb, 1971; Davy, 1979;
Anand et al., 1991) and, possibly, lower in the profile (Grubb, 1971). Small amounts of relict muscovite occur in
duricrust on granitoids. Quartz is abundant (20-25%) over granite, but is much lower (5%) over mafic rocks.
Deeper in the profile, the major change from bauxite to saprolite is marked by the decrease of gibbsite, although in
places gibbsite is found in the lower saprolite at close to the weathering front. Halloysite and/or kaolinite are most
abundant in the saprolite and are the principal weathering products of feldspars and, to a lesser extent, biotite. Feldspars
may weather directly to kaolinite, or via halloysite as an intermediate phase (Anand et al., 1985). Kaolinite crystals
typically have platy morphology and are generally <0.5 μm, whereas halloysite crystals are generally tubular but may
be partly unrolled or platy especially higher in saprolite profiles. Over mafic rocks, smectite clays are abundant in the
saprock and lower saprolite, derived from the weathering of amphiboles and chlorite.
Goethite and hematite generally occur as small platy or sub-rounded crystals in soil and duricrusts. Goethite crystals
range from 160-260 μm, whereas hematite ranges from 180-690 μm (Anand and Gilkes, 1987b; Singh and Gilkes,
1992). Where both Fe oxides occurred in the same sample, hematite crystals were always approximately 60% larger
than goethite crystals. Aluminium substitution for Fe is quite common in both minerals, affecting the unit cell size,
particle size and crystallinity. It is a sensitive indicator of the pedogenic environment of their formation and is
dependent on accompanying minerals (Schwertmann, 1988). Substitution of more than 33 mole % Al does not occur in
either natural or synthetic crystals (Schwertmann, 1985). Aluminum substitution ranges of 22 to 35 mole % in goethites
associated with gibbsite in bauxitic duricrusts in the Darling Range at Boddington, Jarrahdale, Del Park, Huntly and
Willowdale (Anand and Gilkes, 1987b; Anand and Paine, 2002). These values are consistent with those for highly
weathered materials in Australia and elsewhere (Fitzpatrick and Schwertmann, 1982). Separation of morphologically
distinct materials in bauxitic duricrusts in the Darling Range showed that Al substitution in goethite and hematite in the
matrix (mean 25 mole %) and pisolith cutans (mean 25 mole%) was systematically lower than for pisolith cores of (29
mole %) (Anand and Gilkes, 1987a). Fitzpatrick and Schwertmann (1982) also found that goethites in ferruginous
bauxites from South Africa was highly substituted, containing between 20 and 25 mole % Al and was lower in nodule
rinds (14-18 mole %) than cores (20-27 mole %).
There is no significant differences in Al substitution in Fe oxides between lateritic duricrust on granite and dolerite in
the Darling Range (Anand and Gilkes, 1987b), despite the different abundances of Fe, Al and other ions. This implies
that the amount of Al substitution must reflect formation conditions rather than the gross abundance of ions. In
contrast, Singh and Gilkes (1992) found that substitution of Al in goethite was greater for soils on felsic rocks (median
value, 26 mole %) than those on alluvial (17 mole %) and mafic (19 mole %) parent materials. However, no reasons
were given for these differences.
7
GEOCHEMISTRY OF WEATHERING PROFILES
The behaviour of elements during lateritic weathering in the Darling Range regolith has been discussed by several
workers (Sadleir and Gilkes, 1976; Davy, 1979; Davy and El-Ansary, 1986; Ball and Gilkes, 1987; Anand et al., 1991;
Anand, 1994). The variation in element abundances in the profile are a function of bedrock geology and the modifying
influences of weathering. Each element has a particular pattern of enrichment or depletion through the profile. This
can be demonstrated by weathering profiles developed on andesite from Boddington (Figure 7).
The greatest chemical changes between bedrock and the regolith are at the weathering front, the interface between
bedrock and saprolite. Sodium, Ca and Mg are strongly leached from saprolite and the overlying horizons. Potassium
is also lost from alkali feldspars (Anand and Gilkes, 1987c) but is retained throughout much of the profile in muscovite,
which is resistant to weathering. Weathering of biotite to mixed clay minerals and goethite is accompanied by
substantial, but not total, leaching of potassium. Aluminium is enriched in the weathering profile compared to the
bedrock. There is an abrupt rise in Al2O3 at weathering front but, above this, there is little variation throughout the
overlying clay-rich saprolite and mottled clays until the bauxite zone. Iron abundances generally increases upwards
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through the profile although, on granite, there is depletion from the clay-rich upper saprolite. The abundance of Fe 2O3
in the duricrust generally reflect the underlying bedrock. The Fe content of the dolerites (12-20% Fe2O3) is far greater
than that of granite (2-5% Fe2O3). This difference is reflected in the duricrusts, which contain 40-45% Fe 2O3 and 1012% Fe2O3, respectively. Duricrust over andesite contains 25% Fe2O3. Other major changes in Fe content occur where
residual duricrust on crests and upper slopes passes laterally into transported gravels on lower slopes; however, these
changes are not systematic but depend upon the nature of the residual duricrust upslope and on effects related to the
transport of Fe (Hickman et al., 1992). Silica concentrations are also controlled by the nature of the bedrock, mainly
reflecting the abundance of quartz.
Figure 4: Mineralogy of the Darling Range lateritic bauxite profiles. (A) On granitic bedrock : Jarrahdale (modified
from Sadleir and Gilkes (1976). (B) On felsic andesite: Boddington (after Anand, 1994). (C) On metabasalt: Mt
Saddleback bauxite deposits (after Ball and Gilkes, 1985).
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Figure 5: Distribution of poorly crystalline alumina, hematite, maghemite, goethite and gibbsite in the ferruginous zone
of bauxitic weathering profiles from the Boddington gold deposits (after Anand, 1994).
Figure 6. Relationship between Al2O3 and loss on ignition (LOI) in regolith units from Boddington gold deposit (after
Anand, 1994).
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The resistance of quartz to weathering results in its retention throughout the profile into the duricrust, with only small
loss by dissolution. Thus, SiO2 concentration in duricrusts is much greater over the granitoids (15-20% SiO2) than over
dolerite (<5% SiO2). Total silica contents decrease up profile, with a fairly sharp break between the saprolite and the
bauxite zone, due to the alteration of much of the kaolinite to gibbsite.
The distributions of Ti, Zr, V, Cr, Nb, Th and Ga reflect their inertness during weathering, which is due to their relative
chemical immobility (e.g. V, Ti) and/or to the stability of primary and/or secondary host minerals (e.g. Zr in zircon; Ti
in anatase; Cr in chromite). Their abundances all tend to increase upwards through the profile due to the gradual loss of
other components. Chromium, V and Ga tend to accumulate with Fe oxides; Ga is also associated with gibbsite. In
consequence, these elements become concentrated in lateritic duricrust, within which lateral dispersion by colluvial
action can occur during the course of profile evolution.
Of the trace elements, Ba, Sr, Cu, Co, Ni and Zn hosted by feldspars and ferromagnesian minerals are commonly
leached in the saprolite, although a proportion of Co, Cu and Zn is retained in Fe oxides derived from ferromagnesian
minerals. Gold is a significant trace constituent at Boddington, showing moderate concentration in the mid to lower
saprolite and in the duricrust, but with strong depletion in the upper saprolite (Davy and El Ansary, 1986).
There may also be some aeolian accession to the regolith. In the Jarrahdale area, Brimhall et al. (1988) distinguished
two populations of zircons based on their morphology. Euhedral grains were interpreted as having been derived from
the granitic bedrock, whereas rounded grains were regarded as having been derived from aeolian dust and translocated
downwards from the surface by illuviation. Foo (1999) determined the concentrations of Zr in different parts of
pisoliths at Boddington. He found that Zr is greater in the cutans of pisoliths relative to the cores which may support its
derivation in part from external sources during the formation of the cutans.
8
CLASSIFICATION OF LATERITIC DURICRUSTS
8.1
FRAGMENTAL DURICRUST
Fragmental duricrusts generally occur in the lower part of the duricrust. They are composed of angular to subangular,
dark red, hematite-gibbsite fragments of various dimensions ranging from a few mm to more than 30 mm, and
occasionally 100 mm, across, in a pale gibbsite-rich matrix (Figure 8A). The fragments commonly retain remnant
primary lithic fabric with, for example, gibbsite pseudomorphs after feldspars (Figure 8B);voids are infilled with
kaolinite and gibbsite. Similar fabrics in duricrust were observed by Hickman et al. (1992) (Figure 9A,B) and outcrops
having relict bedrock textures are common (Figure 9B).
8.2
FRAGMENTAL-PISOLITIC DURICRUST
Some fragmental duricrusts consist of large fragments containing hematite and gibbsite in a matrix of gibbsite and
goethite that contains pisoliths and ooliths (Figure 8C). In these duricrusts, the fragments are surrounded by red
hematitic and goethite cutans that become paler towards the centre. The pisoliths in the matrix are 2-10 mm in
diameter, with reddish brown to black, hematite-rich cores surrounded by goethite-rich cutans. The ooliths, in
comparison, have cores of goethite, gibbsite and subangular to sub rounded quartz, surrounded by goethitic cutans.
The larger ooliths (<2 mm) are themselves set in a matrix of smaller ooliths (50-100 μm), all cemented by goethite
(Figure 8D).
8.3
PISOLITIC DURICRUST
The pisolitic duricrusts occur at, or close to, the surface and have a greater variety of fabrics and a more complex
mineralogy than the underlying fragmental duricrust. This suggests that individual pisoliths, and different pisolitic
duricrusts, have formed under different weathering environments (Figure 8E and F) (Anand, 1994). Two broad types of
pisoliths have been identified, black and red. Black pisoliths contain minerals that do not occur in the surrounding
matrix. They are massive, 5-15 mm in diameter, and are dominated by maghemite, hematite and PCA. The matrix
between the pisoliths is yellow to light brown, and consists of gibbsite with minor amounts of goethite and hematite. In
contrast, red pisoliths are 2-5 mm in diameter, porous and do not contain maghemite; their mineralogy (hematite,
goethite and gibbsite) is similar to that of the matrix. Some compound pisoliths occur, having earlier generations of
pisoliths within them.
Concentric pisoliths are common. Their nuclei may be massive, composed of either : (a) a mixture of PCA, hematite,
maghemite and gibbsite, or (b) hematite and gibbsite; less commonly, the nuclei are lithic fragments, and are composed
of gibbsite. Most nuclei, irrespective of their composition, are situated at the centre, or slightly offset from the centre of
the pisolith boundaries. Radial desiccation cracks are particularly common within many of the hematite cores,
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suggesting that hematite resulted from dehydration of goethite. Cutans generally comprise at least 30% of the total
pisolith diameter and are mostly concentric, with the innermost cutans having increased thickness so as to fill and
smooth out depressions in the nucleus surface. Discordant boundaries (micro-unconformities) occur between cutans, as
do thin cutans of silt-sized quartz. Where darker and lighter cutans alternate, the lighter cutans are richer in Al, poorer
in Fe and gibbsite-rich.
Figure 7: Box plots showing the distribution of major and minor elements in several regolith materials and felsic
andesite bedrock arranged vertically in order of the regolith stratigraphy.
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Figure 8: Examples of duricrusts. (A) Fragmental duricrust developed on felsic andesite showing hematite-rich
fragments (1) set in a gibbsite-rich matrix (2), Boddington. (B) Photomicrograph of a polished thin section showing
gibbsite pseudomophs after feldspars (1). (C) Fragmental-pisolitic duricrust developed from granite showing large,
irregular fragments (1) containing hematite, quartz and gibbsite in a matrix of gibbsite and goethite that contains
pisoliths (2). The fragments are surrounded by red hematitic (3) and goethitic (4) cutans and become paler towards the
centre. (D) Photomicrograph of a polished thin section showing ooliths of different sizes with cores of goethite
impregnated gibbsite (1) and sub angular to rounded quartz (2) surrounded by goethitic cuatns (3) set in a matrix of
smaller ooliths cemented by goethite (4). (E) Pisolitic duricrust showing pisoliths with cores ranging from earthy red,
hematite-stained gibbsite (1) to dark brown to black hematite and maghemite (2) surrounded by red to brown cutans.
Some compound pisoliths occur (3) having earlier generations of pisoliths within them (4). The matrix is gibbsitic
sandy clay (5) and has numerous small voids (6). (F) Pisolitic duricrust showing concentric pisoliths with a variety of
cores. (G) Massive/vesicular duricrust showing sand cemented by goethite (1) and hematite (2). Harder, more
hematite-rich areas (3) occur along with angular to rounded quartz grains in both the goethite- and hematite-rich matrix.
(H) Photomicrograph of a polished section showing rounded quartz grains (1) embedded in a goethite-rich matrix.
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Figure 9: (A) Lateritic duricrust derived from porphyritic granite. Feldspars crystals have been replaced by gibbsite
pseudomorphs. Original crystal orientation is preserved (after Hickman et al., 1992). (B) Fabric of porphyritic dolerite
preserved in lateritic duricrust (after Hickman et al., 1992).
On lower slopes, some duricrusts have poorly sorted to well sorted pisoliths with yellow goethitic cutans, weakly
cemented by outer cutans. Isolated, round frosted quartz grains, reaching 5 mm in size occur as part of the detrital
framework. Interstices have resulted in a very porous and permeable duricrust with some sandy yellow clay filling.
These duricrusts are formed by erosion of pisoliths from upslope followed by deposition and recementation by goethite
on lower slopes.
8.4
VESICULAR DURICRUST
Vesicular duricrusts occur on the valley floors. These duricrusts (Figures 8G and 8H) commonly contain sandy detritus,
quartz pebbles and a few ferruginous wood fragments, cemented by goethite. Vesicles (1 to 10 mm) in the goethite
matrix vary from rounded to funnel-like or lenticular and irregular. The voids are coated by wart-like protuberances of
yellow ochreous goethite. The duricrusts are formed by the precipitation of Fe oxides in seepages and swamps and are
equivalent to bog iron ore. Where these are organic, as shown by the presence of ferruginized organic matter, goethite
has probably formed from ferrihydrite rather than hematite.
9
DISTRIBUTION OF REGOLITH IN THE LANDSCAPE
Lateritic duricrust occupies gently sloping to horizontal upland areas at 280 to 300 m. Steeper slopes have a thin cover
of transported gravels and bedrock is generally near-surface. There is little duricrust or gravel below 200 m. Blocks of
duricrust, released by headward erosion of streams, degrade to ferruginous gravel on the lower slopes of valleys. On a
local scale, the topography of the weathering front is much more irrregular than that of the top of the duricrust.
Pinnacles of the bedrock and isolated corestones occur high in the weathered profile, locally lessening its thickness.
Contacts between granite and dolerite can be followed in road cuttings and bauxite pit faces to the surface, despite
weathering, indicating that the profile is entirely residual.
Undissected terrain (lateritic upland), characterized by gentle slopes and shallow valleys, has a relatively simple soil
distribution. In general, the soil thickness increases with distance from crests towards the valley floor. Sandy gravel is
predominant on the slopes and crests, whereas sandy soil increases in thickness towards the valley floor. In the
Jarrahdale area, for example, uplands show a toposequence of very gravelly soil (sandy gravel) and lateritic duricrust on
the crests to mid slopes (Figure 10). The crest of the ridge may be occupied by large dolerite dykes represented by
outcrops of fresh rock enclosed in an aureole of lateritic duricrust. The sandy gravels increase in thickness to more than
1 m on the mid slope and lower slope positions, where they overlie various forms of lateritic duricrust and bauxite. The
gravel becomes finer downslope and on lower slopes and is interlayered with sand. ‘Orange earths’ dominate the valley
floor but, in some situations, yellow-brown sand may extend on to them. Podzolic soils and cracking clays are also
prominent in some valley floors. Here, lateritic duricrust is absent and the soils directly overlie saprolite. Pockets of
goethite-rich vesicular duricrust ('bog iron'), formed by chemical or biochemical precipitation of hydromorphically
transported Fe, may occur in valley floors. Outcrops of fresh rock are rare.
On long, steep slopes in dissected terrain, the soil pattern is more complex. Rock outcrops are common, and slump
features, minor landslides and infilled valley-side depressions indicate a prior more dynamic geomorphic environment
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than prevails at present. Orange earths dominate the surface with saprolite an important substrate, although fresh rock
is common. Locally, convex hills are present, and these have greater thickness of soil.
Local variations in lateritic duricrust morphologies occur within a catena or toposequence (i.e. systematic variations
down a slope). Fragmental duricrust occurs on crests and upslope positions, whereas pisolitic duricrust becomes
abundant on mid-slope positions, overlying fragmental duricrust. In some lower slope positions, pisolitic duricrust may
directly overlie saprock or fresh rock. This distribution suggests erosion from crests and upper slopes and deposition on
the lower slopes. Vesicular duricrusts occur in valleys. Similarly, bauxite distribution is related to topographic position,
being best developed on hill slopes, where leaching is greatest, rather than on less freely drained crests or valley floors.
10
REGOLITH-LANDSCAPE EVOLUTION
10.1
EVOLUTION OF WEATHERING PROFILES
The distribution, and mineralogical and chemical compositions, of the lateritic, mostly bauxitic, regolith of the Darling
Range results from the interaction of parent rock, topography, climate, vegetation, drainage and erosion over long
periods of time. Weathering has residually enriched Al and Fe, as a result of Si depletion and the almost total loss of
Mg, Ca, Na and K. These changes are consistent with dissolution of feldspars, amphiboles, biotite, pyroxene and
chlorite, with the retention of Al and some Si in kaolinite and halloysite, of Al in gibbsite and Fe in goethite and
hematite. Similarly, other relatively immobile elements (Cr, V, Ga, Ti and Zr) are concentrated in the upper profile,
especially the duricrust. However, the boundaries of dolerite dykes, quartz pegmatite veins and other primary features
continue without alteration of dips or strike as they pass from bedrock to saprolite and, in places, to the lower duricrust
(Baker, 1976, Sadleir and Gilkes, 1976; Davy, 1979). This, together with the preservation of relict fabrics in the
saprolite, leads to the conclusion that changes from bedrock to saprolite have been nearly isovolumetric. Where
developed, an arenose zone has formed with loss of lithic fabric caused by solution and removal of kaolinite, and
settling of quartz. Progressive weathering under well drained conditions results in increased dissolution of silica and a
bauxite zone is formed.
The characteristics of pisoliths and nodules in duricrusts suggest two modes of formation - in saprock and soil. Nodules
formed in saprock have lithic nuclei that preserve rock fabric. They consist dominantly of gibbsite and hematite;
maghemite is typically absent. Fragmental duricrust has resulted from direct gibbsitization of saprock without
intermediate formation of a kaolinite-rich deep saprolite. Subsequent ferruginization has helped to protect the rock
fabric, except where it is extreme and has progressively destroyed it. Pisoliths formed in soil have homogeneous
nuclei, but may have complex, compound structures; they commonly have concentric cutans. They consist largely of
hematite, maghemite and PCA. Pisolitic duricrusts with the greatest mineralogical diversity have a far more complex
history than mineralogically simpler fragmental duricrusts. Local erosion and deposition would have been continuous
throughout the formation of pisoliths, which are cemented by gibbsite and goethite to form pisolitic duricrust. Pisoliths
with well developed multiple cutans have a complex, cyclic history and have formed by both accretionary and
concretionary processes. Disconformable contacts between cutans, some including thin layers of quartz, indicate that
the cutans formed by accretion. Formation of concentric pisoliths by progressive inward hydration to goethite of a preexisting hematite-rich nucleus, as proposed by Tardy and Nahon (1985), is suggested by irregular dissolution of edges
and scattered remnants of the original nucleus within the cutans. However, the latter situation is uncommon.
The pisoliths have formed by replacement or cementation of soil or colluvium by the solution and migration of Fe and
the precipitation of Fe oxides in the gibbsite matrix or in voids. This was not a single event, but involved multiple
phases of solution and precipitation. The presence of χ-alumina in PCA suggests that aluminium has also been
mobilized and reprecipitated. Transitional aluminas, including χ-alumina, are considered to form by the
dehydroxylation of hydroxides and oxyhydroxides (Misra, 1986; Wefers and Misra, 1987). This can occur by heating,
with the product depending on the rate of heating and the particle size. Thus, dehydroxylation of gibbsite to α–alumina
(corundum) under laboratory conditions proceeds via the χ and k forms whereas dehydroxylation of boehmite proceeds
via the θ, γ and δ forms (Rooksby, 1961). Heating may have been provided naturally by bush fires. However, based on
its morphology, Singh and Gilkes (1995) suggested that the χ-alumina or its precursor was precipitated as a hydrous gel
rather than as discrete crystals. The solubility of Al(OH)3 around pH 4 is fairly high (Misra, 1986) so that significant
amounts of Al may occur in the strongly desilicated, acid soil solution that exists in the near-surface horizons of
bauxitic profiles. In drier seasons, Al may concentrate in soil solution by evaporation and evapo-transpiration, and
precipitate as a hydrous amorphous gel in the soil matrix due to increases in pH and concentration (Hsu, 1989). It is
possible that organic ions in solution would have inhibited the crystallization of gibbsite or boehmite (Hsu, 1989).
Subsequently, this amorphous phase may dehydrate at surface to form transitional alumina and indurate the Fe-rich
matrix in successive seasonal depositions.
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It is unclear when maghemite formed, although it was probably at a late stage, by heating of goethite and hematite by
bush fires (Anand and Gilkes, 1987a). The presence of maghemite thus indicates a very near surface origin, so that the
occurrence of magnetic pisoliths at depths greater than 1-2 m implies later burial or incorporation into soil by
mechanical processes. The materials above buried magnetic pisoliths are thus transported and unrelated to the
underlying regolith or bedrock.
Figure 10: Regolith catena from the Darling Range showing trends in soil and lateritic duricrust following a sequence
from crest to valley floor. Sandy gravels on crest and midslope give way to yellow sandy soils on the lower slopes and
orange earths in the valleys. The gravels become thicker and finer downslope.
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Lateritic duricrusts contain more Fe than the underlying saprolite. Some workers have explained the extra amounts of
Fe in relation to the underlying saprolite by general landscape lowering (e.g., Trendall, 1962; Tardy and Roquin, 1992)
whereas others (e.g., Ollier and Pain, 1996) by lateral and oblique movement of Fe. Trendall (1962), using the Fe
content of granite and the ‘laterite’ over it in Uganda, calculated that 14 m of granite would be required to produce 0.3
m of ‘laterite’ and proposed an overall process of ground surface lowering to account for a surface covered with
‘laterite’. Up to 3 km of surface lowering has been calculated by Tardy and Roquin (1992) for Madagascar. However,
it is not believed that the duricrust in the Darling Ranges is a condensed sequence of this order of magnitude. There
appears to have been little or no lowering of land surface, as shown by the preservation of rock fabrics during the
formation of saprolitic and fragmental duricrust Dolerite dykes and pegmatite veins exposed in the railway cutting at
Jarrahdale can be followed from fresh rock into the duricrust (Sadleir and Gilkes, 1976; Davy, 1979); over the subcrop
of the pegmatite, for example, the duricrust contains coarse, angular quartz. In neither case is there a flattening of the
dip of the contact with the host rock, which would be expected if the duricrust were part of a condensed profile. Terrill
(1956) argued, on similar grounds, that there was no condensed profile over dolerite at Mount Helena and Parkerville.
He pointed out that the Al2O3/Fe2O3 ratio in duricrust over dolerite at Parkerville remained constant from parent rock
through to duricrust. Baker (1976) noted that the occurrence of dolerite dykes was clearly identifiable from the Fe2O3
content of the bauxite horizon, and that the dolerites could be mapped on the basis of the colour and quartz content of
the surface duricrust. However, lowering of a few metres is probable during the formation of pisolitic duricrust. For
example, pisolitic duricrust on felsic andesite has 25% Fe2O3 and the underlying felsic andesite approximately 6%
(Anand, 1994). The concentration factor of 4 is the same for Ti and Zr, from which it can be concluded that the Fe, Ti
and Zr, relatively insoluble elements, were concentrated from the weathering of a greater thickness of underlying rock.
The upper parts of the weathering profiles contain variable amounts of aeolian material, but its source is uncertain. It is
generally difficult to distinguish exotic aeolian constituents from more locally derived alluvial or colluvial material.
10.2
TOPOGRAPHIC CONTROL AND AGE OF WEATHERING PROFILES
Deep weathering profiles at Jarrahdale and Boddington yield late Tertiary palaeomagnetic ages (Dr B. Pillans, ANU,
written communication, July 2002), similar to those reported by Schmidt and Embleton (1976) from the nearby Perth
Basin. In contrast, deep weathering profiles in the eastern Yilgarn yield both early and late Tertiary palaeomagnetic
ages (Table 16 in Anand and Paine, 2002. The absence of older weathering imprints in the western Yilgarn may well be
a consequence of denudation during and since the Mesozoic. The geographical contrast in the depth of post-Permian
dissection from west to east is probably the result of differential uplift associated with breakup of Gondwanaland,
commencing in the early Cretaceous. Minimal uplift to the east allowed preservation of the relict Permian landscape.
There have been various estimates of the amount of material eroded from the Yilgarn Craton (Finkl and Fairbridge,
1979; van de Graaf, 1981; Kohn et al., 1988; Killick, 1998). Finkl and Fairbridge (1979) assumed that there has been
little erosion of the Yilgarn Craton since the Proterozoic, whereas van de Graaf (1981) estimated that up to 400 m of
erosion has been necessary to generate sediments infilling the surrounding basins. Killick (1998) used a similar mass
balance approach to relate the sediment content of the basins surrounding the Western Shield (Pilbara and Yilgarn
Cratons, and intervening Proterozoic Basins), and estimated that that as much as 5.2 km may have been eroded from the
whole Shield. However, he noted that most of the erosion was Mesozoic or earlier, because Tertiary to Recent
sedimentation in the adjacent basins, including the Perth Basin, is dominated by carbonates. Kohn et al., (1998)
obtained similar estimates using apatite fission track thermochronology, which suggested erosion of 3-4 km from the
area of the Darling Range during the late Palaeozoic and, possibly, an equivalent loss during the Mesozoic. There is
strong support for the models of early Phanerozoic denudation along the western margin of the Yilgarn, adjacent to the
Perth Basin. The geomorphological and sedimentary record of Late Palaeozoic glaciation appears to have been
completely removed, with the exception of small outliers of Permian glacial strata and coals preserved in depositional
rifts (Mishra, 1996). It should be noted, however, that extensive denudation models, such as that of van de Graaf
(1981), have been questioned by Sircombe and Freeman (1999) and Cawood and Nemchin (2000) who identified low
concentrations of shield-derived zircon grains in deposits of the Perth Basin. They suggested that this is evidence that
the western part of the shield was not a significant source of sediments during the Phanerozoic and, thus, unlikely to
have experienced extensive stripping.
There has been a great deal of discussion on the role of topography on the formation of bauxitic weathering profiles.
Hickman et al. (1992) summarize many authors’ views on the evolution of bauxite in relation to topography, without
committing to a preferred interpretation. Woolnough (1918) suggested that near sea level peneplain conditions were
required for 'laterite' formation and that its present topographic high positions resulted from subsequent uplift. Playford
(1959) suggested that during lateritization the land surface must have displayed more relief than remains today because
a watertable always shows less relief than its overlying surface. By contrast, Terrill (1948, 1956) implies that the
topography at the time of lateritization closely resembled that seen today. The topography of the bauxitic deposits is
controlled by the nature of the bedrock; faults define stream courses, thicker dolerites underlie ridges (Davy, 1979;
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Hickman et al., 1992). If the bauxite formed in the past on a peneplain, which was then uplifted and subsequently
eroded, the thickest bauxite would not now lie on the mid slopes. If they formed in the past on topography similar to
that of today, it follows that all the erosion that presumably followed the rejuvenation of the streams did not modify the
established bauxite-topography relationships. It appears that the bauxites in the Darling Range are forming today
(Bettenay et al., 1976, Anand and Paine, 2002). This is consistent with the decrease in bauxitic grade with the present
rainfall, i.e. in an easterly direction, and with the work of Dammer et al. (1999) who found the zero age for
cryptomelane from Mundijong and suggested that, at present, intense weathering might still be active in areas with
abundant precipitation (1000 mm/y and more). The dominance of older ages of K-Mn oxides (between 36 and 20 Ma)
obtained in the central part of the Yilgarn Craton, relative to those obtained from southern and southwestern coastal
areas, may reflect the time when climate in the Eastern Goldfields became too arid to enable the formation of K-Mn
oxides in the regolith. This may indicate that differences in climate similar to present conditions between the central
and coastal regions have existed since the late Miocene (Dammer et al., 1999). The seasonal climate provides a
favourable environment for the strong leaching and active formation of gibbsite under the winter rainfall of the present
Mediterranean climate (Anand et al., 1985), compared to lesser leaching and the formation of kaolinite in uplifted
plateaux in savannas having a similar rainfall during the summer (Butt and Zeegers, 1992).
11
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